Forest Ecology and Management 220 (2005) 284–299
www.elsevier.com/locate/foreco
Soil carbon distribution and quality in a montane
rangeland-forest mosaic in northern Utah
Helga Van Miegroet a,*, Janis L. Boettinger b, Michelle A. Baker c,
Julia Nielsen c, Dave Evans a, Alex Stum b
a
Department of Forest, Range, and Wildlife Sciences, Utah State University, 5230 Old Main Logan, UT 84322-5230, USA
b
Department of Plants, Soils, and Biometeorology, Utah State University, Logan, UT 84322-4820, USA
c
Department of Biology, Utah State University, Logan, UT 84322-5305, USA
Abstract
Relatively little is known about soil organic carbon (SOC) dynamics in montane ecosystems of the semi-arid western U.S. or
the stability of current SOC pools under future climate change scenarios. We measured the distribution and quality of SOC in a
mosaic of rangeland-forest vegetation types that occurs under similar climatic conditions on non-calcareous soils at Utah State
University’s T.W. Daniel Experimental Forest in northern Utah: the forest types were aspen [Populus tremuloides] and conifer
(mixture of fir [Abies lasiocarpa] and spruce [Picea engelmannii]); the rangeland types were sagebrush steppe [Artemisia
tridentata], grass-forb meadow, and a meadow-conifer ecotone. Total SOC was calculated from OC concentrations, estimates of
bulk density by texture and rock-free soil volume in five pedons. The SOC quality was expressed in terms of leaching potential
and decomposability. Amount and aromaticity of water-soluble organic carbon (DOC) was determined by water extraction and
specific ultra violet absorbance at 254 nm (SUVA) of leached DOC. Decomposability of SOC and DOC was derived from
laboratory incubation of soil samples and water extracts, respectively.
Although there was little difference in total SOC between soils sampled under different vegetation types, vertical distribution,
and quality of SOC appeared to be influenced by vegetation. Forest soils had a distinct O horizon and higher SOC concentration
in near-surface mineral horizons that declined sharply with depth. Rangeland soils lacked O horizons and SOC concentration
declined more gradually. Quality of SOC under rangelands was more uniform with depth and SOC was less soluble and less
decomposable (i.e., more stable) than under forests. However, DOC in grass-forb meadow soils was less aromatic and more
bioavailable, likely promoting C retention through cycling. The SOC in forest soils was notably more leachable and
decomposable, especially near the soil surface, with stability increasing with soil depth. Across the entire dataset, there
was a weak inverse relationship between the decomposability and the aromaticity of DOC. Our data indicate that despite similar
SOC pools, vegetation type may affect SOC retention capacity under future climate projections by influencing potential SOC
losses via leaching and decomposition.
# 2005 Elsevier B.V. All rights reserved.
Keywords: C sequestration; Decomposition; Leaching; Sagebrush; Forest soils; Soil genesis
* Corresponding author. Tel.: +1 435 797 3175; fax: +1 435 797 3796.
E-mail address: [email protected] (H. Van Miegroet).
0378-1127/$ – see front matter # 2005 Elsevier B.V. All rights reserved.
doi:10.1016/j.foreco.2005.08.017
H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
1. Introduction
Terrestrial ecosystems mitigate increases in atmospheric CO2 via carbon (C) sequestration in vegetation
biomass and soil organic carbon (SOC). While C stored
in aboveground plant biomass and detritus can be
rapidly released following disturbance (e.g., wild fire,
land clearing), C in the mineral soil responds more
slowly to such disturbances (e.g., Johnson and Curtis,
2001; Guo and Gifford, 2002). Temperate, boreal, and
low-latitude forests are significant C sinks (Post et al.,
1982; Jobbagy and Jackson, 2000) followed by
temperate grasslands (e.g., Burke et al., 1989b). Desert
shrub ecosystems sequester less organic carbon (OC)
than forests or grasslands, but a larger proportion of the
OC is stored as SOC (Houghton, 1995). While the
influence of vegetation and climate on C sequestration
in aboveground biomass is well-studied, much less is
known about SOC dynamics (Lal et al., 1997), or about
the C sink strength of montane ecosystems of the
semi-arid West (Schimel et al., 2002), even though
ecosystems such as these occupy >1.6 million km2
globally (Bailey, 1998).
The SOC pool size and distribution reflect past C
inputs and long-term accumulation processes relative to
the cumulative hydrologic and gaseous C losses from
the soil. Vegetation type and productivity influence C
input rates and distribution in the soil. In forests, <50%
of net primary productivity (NPP) is allocated to
roots (Coleman, 1976; Swift et al., 1979; Raich and
Nadelhoffer, 1989), and OC additions to the soil occur
chiefly via litterfall to the soil surface resulting in the
formation of distinct O horizons. In contrast, most C
allocation in shrubs and grasses is to roots (Coleman,
1976; Swift et al., 1979; Tiedeman, 1987), belowground
C storage comprises most of the ecosystem OC (e.g.,
Prichard et al., 2000), and OC addition occurs mainly
via fine root turnover in the near-surface mineral soil
horizons (Swift et al., 1979; Knight, 1991), resulting in
limited surficial litter accumulation. Differences in
vegetation type and the mode of C input to the soil
further impact SOC quality (e.g., Barth and Klemmedson, 1982; Johnson, 1995; Quideau et al., 2001). Such
differences in SOC distribution and dynamics among
vegetation types are often reflected in divergent soil
forming processes and horizon properties (e.g., Foth,
1984; Quideau and Bockheim, 1996). Conifer forests
generate litter that is relatively slow to decompose, but
285
tends to result in the production of highly soluble
organic acids (Hongve et al., 2000), mostly fulvic acids
(Kononova, 1966 as cited in Knight, 1991; Stevenson,
1982; Quideau and Bockheim, 1996), involved in the
podzolization process. The fine root turnover of grasses
tends to produce less-soluble humic acids and more
stable humins (Stevenson, 1982; Anderson and Coleman, 1985), responsible for the formation of thick and
distinct A horizons and Mollisols under prairies (e.g.,
Foth, 1984).
It is not clear to what extent the above patterns in
SOC distribution and dynamics among broad vegetation types also persist in seasonally dry ecosystems,
where NPP, decomposition and hydrologic fluxes are
limited by water availability (Knight, 1991). There is
evidence that changes in temperature and/or moisture
or disturbances may induce SOC quality changes,
even for a given vegetation type or soil substrate
(Homann and Grigal, 1992; Amelung et al., 1997),
which in turn may influence SOC storage (Jackson
et al., 2002), SOC stability, or potential for C losses via
leaching or decomposition (Anderson and Coleman,
1985; Zech et al., 1989; Homann and Grigal, 1992;
Christ and David, 1996). On a global scale, climatic
control of C dynamics is thought to lead to equilibrium
SOC levels along broad temperature and precipitation
gradients (e.g., Post et al., 1982; Kirschbaum, 1995).
This would imply that vegetation type has limited
impact on C sequestration potential of soils within a
certain climatic region.
Seasonally dry montane ecosystems of the Intermountain West (IMW) are typified by structurally
heterogeneous forest-rangeland mosaics. Yet, relatively little is known about SOC quality and dynamics
in semi-arid systems or their variability across the
landscape. To that end, we initiated a pilot study in
Utah State University’s T.W. Daniel Experimental
Forest (TWDEF) in northern Utah, where a mosaic of
forest (conifer and aspen) and rangeland (grass-forb
and sagebrush) ecosystems typical of the IMW,
occurs on non-calcareous soils under similar climatic
conditions. The objective was to determine whether
meaningful differences existed in the quantity,
distribution, and quality of SOC in soils found under
different vegetation types. Such data could provide
valuable information on potential variability in C
storage and retention in semi-arid environments, and
the stability of the SOC pools under changing
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H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
climate. To represent a longer term C sink, we
focused our investigation mainly on the OC in
mineral soil horizons, which are generally less
responsive to disturbances (Johnson and Curtis,
2001; Guo and Gifford, 2002). Furthermore, in
defining SOC quality, we focused on function rather
than operational fractionations, and specifically
considered decomposition and leaching potential of
SOC. These characteristics are critical to stability and
retention of SOC as they will influence C losses as
CO2 via microbial decomposition and as dissolved
organic C (DOC) via leaching.
2. Materials and methods
2.1. Study site
The TWDEF is located at an elevation of 2600 m,
approximately 30 km NE of Logan, Utah (41.868N,
111.508W). Average annual precipitation is 950 mm,
80% of which accumulates as snow. Snowmelt
typically occurs from mid-May to mid-June. Monthly
rainfall is low between May and October, with lowest
monthly precipitation (<2 cm) typically occurring in
July. Average low temperature is around 10 8C in
January; highest mean monthly temperature (14.5 8C)
occurs in July (Schimpf et al., 1980; Skujins and
Klubek, 1982). Summer grazing by cattle and sheep
has occurred since the late 1800s (Schimpf et al.,
1980) but was greatly reduced coincident with fire
suppression since 1910 (Wadleigh and Jenkins, 1996).
There is ample evidence in the literature that unless
areas are severely overgrazed, few consistent changes
in SOC occur (Manley et al., 1995; Bruce et al., 1999).
Hence, we do not consider past grazing history a
significant driver of SOC accumulation in our study
area. Following an increase in fire frequency during
the 1856–1909 settlement period, fire frequencies
have significantly declined, and there is no evidence of
fire in the area since 1910 (Wadleigh and Jenkins,
1996).
Our study was conducted in a 150 ha area with a
mosaic of vegetation communities, all in close
proximity to each other and characterized by similar
elevation, aspect, climate, geomorphology, and geology (Fig. 1). Forested communities include aspen
forest [Populus tremuloides Michx] and conifer forest,
predominantly Engelmann spruce [Picea engelmannii
Parry] and subalpine fir [Abies lasiocarpa (Hook)
Fig. 1. Air photo of the T.W. Daniel Experimental Forest showing locations of the pedons [A, Aspen; AC, Mixed Aspen-Conifer; C, Conifer; G,
Grass-Forb; GC, Grass-Conifer ecotone; S, Sagebrush]. Solid lines indicate boundaries of the aspen forest; dotted lines surround sagebrush
steppe.
H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
Nutt], and lodgepole pine [Pinus contorta] stands.
Non-forested communities include open meadows
consisting of a mixture of grasses and forbs, and areas
dominated by sagebrush [Artemisia tridentata].
The soils in the study area are carbonate-free and
generally well drained, formed in eolian deposits
overlying residuum and colluvium from the Wasatch
formation (tertiary: middle and lower Eocene)
dominated by roughly stratified, poorly sorted
conglomerate a few hundred meters thick (Dover,
1995). This allowed us to focus on SOC dynamics
without the interference of inorganic C (carbonates).
2.2. Sampling and analysis
A soil survey of TWEDF conducted by Utah State
University’s Soil Genesis, Morphology, and Classification course started in September 2000 and is
ongoing. Soil pedons representative of major vegetation types and landscape components were exposed by
manual excavation. Soil morphology was described
following standards methods (Soil Survey Division
Staff, 1993), including horizon depth, color, texture,
structure, and volume of rock fragments. Six pedons
were described in September 2000 under mixed
spruce-fir conifer (C-0a, C-0b), aspen (A-0), mixed
aspen-conifer (AC-0), and grass-forb meadow
(G-0a, G0-b). In September 2001, four pedons were
described under conifer (C-1), aspen (A-1), mixed
aspen-conifer (AC-1), and sagebrush (S-1), and the
near-surface mineral horizons were sampled for
further analysis. In September 2002, five pedons
under conifer (C-2), aspen (A-2), grass-forb meadow
(G-2), sagebrush (S-2), and the shaded ecotone
between grass-forb meadow and conifer (GC-2) were
described and sampled by genetic horizon to 129 cm
soil depth (Fig. 1).
Genetic horizon samples from 2002 were air-dried
and sieved <2 mm, and selected chemical and
physical analyses were performed. Exchangeable
base cations were extracted with 1 M NH4OAC at
pH 7.0 using a vacuum extractor (Soil Survey
Laboratory Staff, 1996) and determined using and
inductively coupled plasma spectrometer (ICP) (Iris
Advantage, Thermo Electron, Madison, WI). Particle
size distribution (PSD) was determined by wet sieving
and pipette analysis (Gee and Bauder, 1986) after
removal of organic matter with 5% NaOCl adjusted to
287
pH 9.5 with HCl. Total C (assumed to equal SOC in
non-calcareous soils) and total N concentrations were
determined by dry combustion using a CHN analyzer
(Leco CHN 1000, Leco Corp., St. Joseph, MI). Bulk
density of the <2 mm fraction was estimated according
to texture (1.2 Mg m3 silt loam surface horizon;
1.3 Mg m3 silt loam subsurface horizon, loam;
1.4 Mg m3 clay loam, silty clay loam; 1.5 Mg m3
sandy loam; 1.6 Mg m3 loamy sand). The total C
pools for each genetic horizon were calculated by
multiplying C concentration, bulk density, horizon
thickness, and the volume fraction of rock-fragment
free soil by a unit conversion factor. Total C for the
whole pedon was calculated by summing the C content
for all horizons to the depth of sampling.
The soil organic matter (SOM) content was
measured as mass loss-on-ignition (Nelson and
Sommers, 1996) on a representative subsample from
each genetic soil horizon sampled in 2002 and the
near-surface horizons sampled in 2001. Water-soluble
organic C, which in this paper is further referred to as
dissolved organic C, was extracted from the same
sample set with deionized water using a 2:1 (water to
soil) ratio by shaking at room temperature (Baker
et al., 2000). The extracts were filtered using 0.22 mm
filters and DOC concentration was measured using
wet persulfate oxidation (Menzel and Vaccaro, 1964)
on a TOC analyzer (Oceanographic International,
College Station, TX). Decomposable or bioavailable
DOC (BDOC) was determined by inoculating a
subsample of extracted DOC with microorganisms
back-flushed from filters used to prepare DOC
samples and incubating sample in the dark for 30
days (Servais et al., 1987; Baker et al., 2000).
Following incubations, the remaining DOC content
was measured using wet persulfate oxidation as
described above, and BDOC was calculated as the
difference between initial and final DOC concentration. DOC quality in terms of aromaticity was
determined as specific ultra violet absorbance (SUVA)
at 254 nm (Westerhoff and Anning, 2000).
The relative decomposability of SOC was determined from long-term aerobic laboratory incubation
(Paul et al., 2001) of mineral soil samples taken in
2002 from the three upper mineral horizons in the five
pedons. Approximately 80 g of field-moist, unsieved
soil was incubated aerobically at 25 8C in glass jars
(two laboratory replicates per sample). Two blanks
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H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
(incubation jars without soil) were included in the
design. Incubation jars were aerated weekly and
soils were periodically weighed and water added to
maintain initial soil moisture contents. CO2 evolution
was measured periodically (initially, weekly, or
biweekly for the first 8 weeks and monthly thereafter)
using sodalime as a CO2 trapping agent. A correction
factor of 1.41 was applied to weight differences to
account for water loss during drying of sodalime at
105 8C (Edwards, 1982; Zibilske, 1994). For simplicity sake, the pedons will be identified by their
vegetation cover throughout the rest of the paper.
3. Results and discussion
3.1. Soil morphology and genesis
As expected, forest soils had characteristic O
horizons (Table 1). These were thickest and most
pronounced where conifers were present [6 cm under
spruce-fir forest (C-0a, C-0b, C-1, C-2), 5 cm under
mixed conifer-aspen forest (AC-0, AC-1)] (Fig. 1). The
O horizons were generally thinner (3 cm) in soils
under aspen (A-0, A-1, A-2), reflecting the more easily
decomposable foliage (e.g., Stump and Binkley, 1993;
Prescott et al., 2000). Rangeland soils, i.e., soils under
grass-forb (G-0a, G-0b, G-2, GC-2) and sagebrush
cover (S-1, S-2), lacked O horizons. All soils had A
horizons. Surprisingly, there was little difference in the
thickness of A horizons (A plus AB transition horizons)
between soils of rangeland (meadow, sagebrush) and
conifer ecosystems. The A horizon thickness ranged
from 31 to 35 cm under sagebrush, 12 to 36 cm under
conifers, 17 to 29 cm under grass-forb meadows, and
was about 19 cm in the ecotone between conifer forest
and meadow. Soils under aspen had a slightly thinner A
horizons (6–15 cm under aspen; 6–8 cm under mixed
aspen-conifer). The color of A horizons in most soils
were similar (dry 10YR to 7.5YR hue, 4–6 value),
except in one grass-forb meadow soil (G-0b) located in
a slight water- and sediment-gathering concavity which
was darker (dry 2.5YR to 5YR hue, 3 value).
Soils under forest vegetation or shade tended to
have thicker E and EB or BE transition horizons,
indicating more leaching. This is consistent with
previous research in two subalpine grass-forb meadows in the TWDEF which showed that upper
horizons retained moisture longer following spring
snowmelt in soil shaded by spruce-fir tree islands
compared to exposed meadow soils (Van Miegroet
et al., 2000). Two aspen pedons on steep slopes (A-0,
A-1; 22–27% slopes) lacked E horizons, likely caused
by higher rates of runoff and erosion. Most rangeland
soils lacked E horizons; greater belowground biomass
and fine-root turnover coupled with less leaching
resulted in the junction of A and B horizons.
All soils demonstrated an accumulation of clay in
the subsoil and had B or B transition horizons (Bt,
BEt) most of which classified as argillic horizons (Soil
Survey Staff, 1999). Base cations were also high in
the subsoil (data not shown). Thus, most soils were
Alfisols. Some soils under aspen and grass-forbs that
had dark surface horizons at least 18-cm deep and
pH about 6 or higher throughout their depth were
Mollisols. The temperature regime of most soils was
classified as cryic (4 8C mean annual soil temperature and relatively cool summer soil temperatures) and
the moisture regime as udic (moist for most of the
year). Soils under sagebrush were slightly warmer
(frigid temperature regime: 5 8C mean annual soil
temperature, relatively warm summer soil temperature) and drier (xeric moisture regime: dry in the upper
part for at least 45 consecutive days after the summer
solstice).
Limited water movement plays a significant role
in soil development in this environment. Soils under
conifer forests did not develop into Spodosols
primarily because of lack of sufficient water flux
through the soil during a significant part the year.
Alfisols and Mollisols are more typical of forest and
rangeland soils seasonally dry environments; their
high base status demonstrates that nutrient leaching
losses are low due to moisture restrictions.
3.2. SOC pools and distribution in the mineral
soil
Both soil organic matter (data not shown) and OC
concentrations were highest in the upper part of the
mineral soil and decreased with depth (Table 1). The
OC concentration in the upper A horizon near the
soil surface was higher in the forest pedons (35–
68 mg g1) and lowest in the sagebrush pedon
(16 mg g1). The OC concentration declined sharply
with depth in the forest soils, especially under aspen,
H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
289
Table 1
Morphological, physical, and chemical properties of soil pedons sampled in 2002
Field
pH
Field
texture
Clay
(%)
Volume
2-mm (%)
Estimated bulk
density (Mg m3)
Conifer (C-2): Loamy-skeletal, mixed, superactive Umbric Haplocryalf
Oi
0–5.5
nd
nd
nd
nd
A (with O) 5.5–7
10YR 3/2
67.6
3.35
24
A1
7–17
10YR 5/3
15.6
0.66
28
A2
17–42
10YR 5/3
13.2
0.61
25
E
42–61
10YR 6/3
5.6
0.12
53
EB
61–76
7.5YR 7/4 1.9
<0.10
–
Bt1
76–112
7.5YR 6/4 2.0
<0.10
–
Bt2
112–150
7.5YR 6/4 2.2
<0.10
–
nd
5.6
5.4
5.4
4.8
4.8
4.6
4.6
na
gr L
gr L
gr L
gr L
gr LS
vst SL
vgr SL
na
15
15
13
12
10
13
16
na
0
20
15
17
30
50
50
na
1.2
1.3
1.3
1.3
1.5
1.3
1.5
Aspen (A-2): Fine, smectitic
Oi
0–1
A1
1–4
A2
4–16
E1
16–33
E2
33–54
BEt
54–68
Bt
68–98
E’
98–102
BE’t
102–130
nd
6.0
6.2
5.4
5.2
5.1
5.1
4.9
4.8
na
gr L
gr SiL
vgr L
xcb SiL
SCL
C
SL
CL
na
14
13
11
10
28
52
18
36
nd
15
15
55
70
5
0
0
0
na
1.2
1.3
1.3
1.3
1.4
1.5
1.3
1.4
Meadow–conifer ecotone (GC-2): Fine-loamy, mixed, superactive Oxyaquic Haplocryalf
A1
0–13.5
10YR 5/4
19.7
0.98
23
5.4
gr L
A2
13.5–18.5 10YR 5/3
16.5
0.86
22
5.4
L
E
18.5–38.5 10YR 7/3
8.7
0.40
26
5.6
vst L
EBt
38.5–53
10YR 7/3
3.0
0.11
33
5.8
SCL
Bt
53–78
10YR 6/3
3.5
0.17
24
5.5
st CL
Bt/E1
78–94
7.5YR 4/4 3.1
0.13
27
5.4
CL
Bt/E2
94–110
7.5YR 4/4 3.4
<0.10
–
5.3
CL
E/Bt
110–136
7.5YR 4/4 2.3
0.10
26
5.3
gr SCL
B’t
136–160
5YR 5/6
1.4
<0.10
–
4.8
SC
19
18
19
21
33
35
30
30
43
20
5
55
10
20
5
5
15
0
1.2
1.3
1.3
1.3
1.4
1.4
1.4
1.4
1.5
Grass forb meadow (G-2): Fine, smectitic
A1
0–11
10YR 4/4
A2
11–18
10YR 4/3
A3
18–28
10YR 4/3
E
28–50
10YR 5/3
EBt
50–73
10YR 7/4
Bt1
73–92
2.5YR 5/6
Bt2
92–111
2.5YR 5/8
Bt3
111–145
10R 5/8
Horizon
Depth
(cm)
Color
(dry)
C
(mg g1)
Typic Haplocryalf
nd
nd
10YR 6/3
34.6
10YR 6/3
13.8
10YR 7/3
5.3
10YR 7/3
1.8
2.5YR 4/6 2.7
10R 4/6
1.7
7.5YR 8/4 1.1
2.5YR 6/6 1.0
N
(mg g1)
nd
1.95
0.85
0.34
<0.10
0.19
<0.10
<0.10
0.11
Molar
C/N
nd
21
19
18
–
17
–
–
10
Umbric Palecryalf
23.4
1.40
20.5
1.30
11.0
0.87
9.2
0.69
2.9
<0.10
1.7
<0.10
1.4
<0.10
1.0
<0.10
20
18
15
16
–
–
–
–
5.0
5.0
5.3
5.1
5.0
4.6
4.5
4.5
SL
SL
SCL
SCL
cb SCL
C
SC
C
26
28
29
36
23
43
40
43
5
5
12
7
20
0
0
0
1.2
1.4
1.4
1.4
1.3
1.5
1.4
1.5
Sagebrush (S-2): Fine, smectitic, frigid Ultic Haploxeralf
A1
0–16
7.5YR 5/3 15.8
0.99
A2
16–31
7.5YR 5/3 13.6
0.77
EBt/Bt
31–42
7.5YR 5/4 8.5
0.42
31–42
5YR 5/4
4.8
0.22
Bt1
42–61
5YR 4/6
3.2
0.10
Bt2
61–82
5YR 4/6
1.7
<0.10
Bt3
82–129
2.5YR 4/6 1.3
<0.10
19
20
26
23
37
–
–
5.4
5.4
5.3
5.1
5.1
4.7
4.7
SiL
SiL
CL
C
CL
C
C
22
22
42
33
36
49
52
15
25
21
–
0
0
0
1.2
1.3
1.5
1.4
1.4
1.5
1.5
Abbreviations: nd: not determined, na: not applicable, v: very, x : extremely, gr: gravelly, cb: cobbly, st: stony, bd: bouldery, LS: loamy sand, LFS:
loamy fine sand, SL: sandy loam, FSL: fine sandy loam, COSL: coarse sandy loam, SiL: silt loam, L: loam, CL: clay loam, SCL: sandy clay loam,
SiCL: silty clay loam, C: clay, SC: sandy clay, and SiC: silty clay.
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H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
where the soil OC concentration dropped to 5 mg g1
below 16 cm soil depth. The OC concentrations were
similar in soils under the grass-forb cover, irrespective
of location in the open meadow or the forest-meadow
ecotone, and remained elevated to a greater depth
(40–50 cm). Below 60 cm, OC concentrations did not
differ much among pedons and typically ranged
between 1 and 3 mg g1 (Table 1).
The SOC concentrations in the conifer, aspen, and
grass-forb meadow soils are similar to values reported
earlier for the same study site (Schimpf et al., 1980;
Van Miegroet et al., 2000). They are higher than for
comparable horizons in sandy pine and prairie soils in
southeastern Wisconsin (Quideau and Bockheim,
1996), but this may be partly due to the higher
percent sand in soils at that study site. The SOC
concentrations and their vertical distribution in grassforb meadow soils are within the range observed under
native prairie vegetation in the Great Plains (Amelung
et al., 1997). The A horizon concentrations are lower
than in Mollisols and native grassland soils, but similar
to cultivated grassland soils in Montana (Sims and
Nielsen, 1986) and higher than beneath bluebunch
wheatgrass in eastern Washington (Bolton et al., 1990,
1993). The SOC concentrations in the aspen pedon are
similar to values for aspen soils in interior Alaska in
reported by O’Neill et al. (2002). The A horizon
(0–16 cm) SOC concentrations are significantly
lower than measured in closed-canopy stands at
similar elevation in the Rocky Mountains of Colorado
(Giardina et al., 2001), but comparable to Ae
concentrations reported for aspen in Saskachewan
(Huang and Schoenau, 1996). The SOC concentrations under sagebrush are on the low end of values
measured in upper 0–20 cm of sagebrush soils in
eastern Oregon (Doescher et al., 1984), in eastern
Washington (Bolton et al., 1993), and at various soil
depths in Wyoming (Burke et al., 1987; Burke, 1989),
but similar to those reported for Utah (Charley and
West, 1975).
Estimated total SOC in this forest-meadow mosaic
ranged from a low of 50 Mg ha1 (0–130 cm) in the
aspen pedon to a high of 109 Mg ha1 (0–145 cm) in
the grass-forb pedon (Fig. 2). The total SOC was
similar under conifer, sagebrush and in the meadowforest ecotone, around 85 Mg ha1. In all cases, more
than 50% of the total mineral soil C pool was
contained in the A horizons (upper 30–40 cm). The
aspen pedon was somewhat of an outlier in that had a
lower SOC pool, half of which was stored in the upper
16 cm of the mineral soil. Soil N was primarily
concentrated in the upper 50–60 cm of the mineral soil
(Table 1), following patterns in SOC accumulation.
The C/N ratio fluctuated around 20 (range: 15–28) in
the upper soil horizons and did not differ among
pedons.
Our calculated total SOC values fall within the
range of worldwide averages for desert shrub,
sclerophyllous scrubland, and temperate grasslands
and woodlands (Lal et al., 1997; Jobbagy and Jackson,
2000), and those reported for forest-prairie ecotones in
Missouri (Hammer et al., 1995). The SOC contained
in the upper 50 cm of grass-forb and meadow-conifer
ecotone soils and in the A horizon of the conifer soil is
on the high end of values reported for Wisconsin
prairie and pine soils, respectively (Quideau and
Bockheim, 1996). Our SOC estimate for the conifer
pedon (0–150 cm) is midway between values for the
upper 100 cm of Alfisols and Spodosols in Denmark
(Verje et al., 2003). It is within the range reported in
the Great Lakes States for pines, but below those for
balsam fir, and intermediate between black and white
spruce in Alaska (O’Neill et al., 2002). Our aspen
pedon estimate, however, is considerably lower than
published values (Grigal and Ohman, 1992; Rollinger
et al., 1997; O’Neill et al., 2002), and could be the
result of the higher gravel content in the upper 50 cm
(depth-weighted about 48% compared to 10–30% for
the other pedons). Total SOC in the upper 15 cm under
sagebrush falls in the middle of values reported by
Burke (1989) for sites in Wyoming. The lower SOC
accumulations in forest pedons at our site may be due
to water limitations on NPP (Knight, 1991) and/or
lower stand densities. Also, the reported SOC pools
here are our best approximations based on estimated
bulk density and coarse fragment volume, which are
known to contribute to significant variation in total
SOC (Hammer et al., 1995; Homann et al., 1995).
3.3. SOC quality
Are these SOC pools similar in stability and
behavior, or could some of the observed differences in
C accumulation be tied back to distinct SOC quality
characteristics? Are SOC quality differences important to stability and retention of SOC in the different
H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
291
Fig. 2. Total SOC (Mg ha1) by genetic horizon in the (a) conifer, (b) aspen, (c) meadow-forest ecotone, (d) grass-forb meadow, and (e)
sagebrush pedon.
ecosystems? We specifically considered indicators of
potential C losses as CO2 via microbial decomposition
and as DOC via leaching.
Water-extractable OC levels (DOC) used as an
indicator for SOC solubility and leaching potential,
were significantly different among pedons ( p = 0.004;
log–log transformated to account for large and
unequal within-pedon variance). The highest DOC
concentrations were measured in the upper horizons of
the forest soils, consistent with higher total OC
concentrations close to the surface of forest soils.
Conifer DOC levels averaged 0.03 mg g1, similar to
those reported for Douglas-fir in Oregon (Jandl and
Sollins, 1997). The aspen pedon consistently yielded
higher DOC throughout the soil profile in 2002
(Fig. 3). The DOC was 2.5 mg g1 in the upper few
cm, decreased to 0.03 mg g1 in the E horizon (24–
60 cm), then increased again below 60 cm depth. This
pattern possibly reflects downward transport of DOC
from the upper SOM-rich horizons and relative
accumulation, associated with a decline in hydraulic
conductivity at the interface with the argillic horizon
292
H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
Fig. 3. Water-extractable DOC concentrations (mg g1 dry soil) in
the different soil pedons. Open symbols refer to 2001 near-surface
soil samples; solid symbols to 2002 soil samples; DOC depth
profiles are linked by lines [meadow, grass-forb meadow; ecotone,
meadow-forest ecotone].
and/or DOC precipitation to the clay surfaces (e.g.,
Schoenau and Bettany, 1987; Quideau and Bockheim,
1996). The DOC levels in the upper A (0–4 cm),
which likely contained some O horizon material, are
similar to water-soluble SOC concentrations (considered equivalent to DOC in our study) measured in
September-October in the organic F (1.2–3.1 mg g1)
and H (0.9–1.9 mg g1) layers under aspen forests in
Saskatchewan (Huang and Schoenau, 1996). That
study also showed a sharp decline in DOC with depth,
concentrations around 0.06–0.08 mg g1 in the Ae,
and the highest DOC in October. Such a spatial and
temporal pattern reflects the mobilization of soluble
OC as the first step in the decomposition of fresh litter
commonly observed in forests (Schoenau and Bettany,
1987; Qualls et al., 1991; Homann and Grigal, 1992;
Hongve et al., 2000). The lower DOC in the conifer
pedon compared to the aspen pedon is consistent with
the findings by Hongve et al. (2000), indicating that
deciduous litter gives off proportionally more DOC
than conifers under similar climate. The lowest DOC
levels were measured in the upper 60 cm of grass-forb
meadow and meadow-forest ecotone soils, while
DOC concentrations in surficial sagebrush soils were
intermediate. The differences in extractable DOC
between forest and grass-forb meadow soils in our
study are consistent with lysimeter data showing lower
solubility of OC under grasses (e.g., Quideau and
Bockheim, 1996).
The DOC was poorly correlated with SOM
(R2 = 0.2; p = 0.001) and uncorrelated with SOC
concentrations (R2 = 0.01; p = 0.55) suggesting that
differences in DOC levels were not driven by OC
quantity and distribution, but rather indicated inherent
differences in SOC quality, notably differences in
solubility, and associated differences in downward
transport. Across the entire dataset (excluding the one
aspen soil outlier with a DOC value of 2.5 mg g1;
Fig. 3), we found a significant positive linear
relationship between the amount of labile (BDOC)
and total DOC (Fig. 4), similar to observations in
agricultural soils by Zslonay and Steindl (1991).
However, there were distinct differences in the
decomposability of the DOC among forest and
rangeland soils: The DOC produced in soils under
grass-forb cover was highly labile, as 100% DOC of in
grass-forb meadow soils and 76–100% of DOC (mean:
89 9%) in meadow-conifer ecotone soils disappeared within a 30-day period during laboratory
incubations. Horizons from the sagebrush pedon fell
slightly below the 1:1 line, indicating slightly lower
DOC decomposability (mean: 76 17%) compared
Fig. 4. Relationship between DOC concentrations (mg g1 dry soil)
and labile DOC (amount of DOC consumed during 30-day laboratory incubation in mg g1 dry soil) in soil samples collected under
the different vegetation types. Open symbols refer to 2001 nearsurface soil samples; solid symbols to 2002 soil samples [meadow,
grass-forb meadow; ecotone, meadow-forest ecotone].
H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
to the grass-forb rangeland type. In contrast, the higher
DOC levels extracted from the forest soils were also
more stable compared to those from the rangeland
soils, with % decomposability ranging between 34 and
87% under conifer (mean: 62 19%) and between 28
and 100% under aspen (mean: 60 24%).
The DOC decomposition rates are in the general
range of those noted for agricultural soils (Zslonay and
Steindl, 1991). Our DOC decomposability values for
forest soils are similar to those measured by Hongve
et al. (2000) for various deciduous species during a 60day incubation (60–70%), but higher than for spruce
(nearly 0), or the 10–40% reported in the literature for
pines and various hardwood species (Qualls and
Haines, 1992; Yano et al., 1998; Corre et al., 1999).
Our estimates for the decomposability of DOC in
grass-forb soils are also higher than the 40% and 50%
reported for cool-season and warm-season grasses,
respectively (Corre et al., 1999).
The SUVA readings expressed per mg DOC
extracted showed considerable variability within each
pedon (CV: 43–58%). They were highest in the conifer
soils (mean SUVA absorbance 100 mg1 DOC:
3.9 1.3), and lowest in the grass-forb meadow
(mean: 1.4 0.7) and meadow-forest ecotone soils
(mean: 1.3 0.6), indicating that the DOC produced
under conifers was relatively more aromatic compared
to that produced under grasses and forbs. The
aromaticity of DOC extracted from aspen (SUVA
mean: 1.8 1.0) and sagebrush soils (SUVA mean:
1.9 0.9) was intermediate between that of conifer
and grass-forb soils (Fig. 5). Hongve et al. (2000)
similarly showed greater SUVA values for spruce
compared to percolates from deciduous litter. Across
all samples, the DOC decomposability and aromaticity as expressed by SUVA readings were inversely
correlated (N = 46; R2 = 0.26; p < 0.001). This suggests a connection between C structure and behavior,
with a higher abundance of aromatic groups contributing to greater resistance of DOC to microbial
decomposition, a conclusion also supported by Qualls
and Haines (1992) and Hongve et al. (2000). Our data
further suggest that low DOC levels in soils under
grass-forb cover may, in part, be due to the rapid
turnover of the soluble C by heterotrophic decomposers. The latter would lower the potential hydrologic C losses and, in effect, promote retention of C in
the upper soil by continuously cycling the OC between
293
Fig. 5. Relationship DOC decomposability (%DOC consumed
during 30-day laboratory incubation) and DOC aromaticity (SUVA
absorbance 100 per mg DOC) in soil samples collected under the
different vegetation types. Open symbols refer to 2001 near-surface
soil samples; solid symbols to 2002 soil samples [meadow, grassforb meadow; ecotone, meadow-forest ecotone].
the solid (microbial biomass) and the soluble phase.
Microbial assimilation of DOC entails C losses as
CO2, and the extent to which a decline in hydrologic
losses (DOC leaching) is countered by an increase in
gaseous losses (DOC decomposition) depends on the
microbial C use efficiency and its controls. Because
we did not measure DOC production rates in this
study, we could not evaluate the role of differential
DOC production rates in causing differences in DOC
levels in this forest-rangeland mosaic. In modeling the
controls of DOC flux and SOC content of forest soils,
Neff and Asner (2001) showed soils with highly
aromatic or hydrophobic DOC were less likely to
leach given similar water availability. However, no
correlation between DOC extracted and aromaticity
emerged from our dataset (R2 = 0.01; p = 0.42).
The cumulative CO2 release over the 10-month
incubation was generally highest in the upper soil
samples, especially in the forest soils (Fig. 6a). As was
the case for total and soluble OC, forest soils exhibited
a sharp change in SOC turnover with increasing depth
below the soil surface. Differences in cumulative CO2
production between the upper three horizons were
generally less pronounced in the rangeland soils. Daily
CO2 evolution rates normalized for OC content
indicated qualitative differences of the SOC across
the landscape (Fig. 7). The initial amounts of CO2
294
H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
Fig. 6. Total amount of CO2 released (mg CO2 g1 soil) (a) and
total CO2 release per gram C (b) during 10 months of aerobic
incubation of upper three mineral soil horizons from the five soil
pedons.
released per g of OC were highest in the upper two
mineral horizons (5–17 cm) of the conifer soil and the
upper 4 cm under aspen. They declined exponentially
with time, and by the end of the incubation period, CO2
evolution rates had dropped to around 1 mg CO2 g1 C
24 h1. The cumulative CO2 release per unit C was
quite comparable for upper forest soil horizons, ranging
between 500 and 550 mg g1 C over 10 months and
representing a little over 13% loss of SOC through
decomposition (Fig. 6b). These cumulative and
temporal trends were similar to those reported by
Giardina et al. (2001) for the upper 15 cm of pine and
aspen soils in The Rocky Mountains. The CO2 release
rates from grass-forb meadow and meadow-forest
ecotone soils were generally lower throughout the
incubation period (mean: <2 mg CO2 g1 C 24 h1),
did not decline sharply with time, and did not differ
among three soil horizons. The observed respiration
kinetics suggests a higher content of labile C in the
upper forest soils (Zak et al., 1993). Ross et al. (1999)
similarly observed a sharp decline in respiration with
increasing depth in pine soils, although their rates were
substantially higher than those reported here. In another
study they also measured greater CO2 release only in
upper forest soils compared to grasslands in montane
sites in New Zealand, with differences between
vegetation type diminishing below 10 cm (Ross
et al., 1996).
Sagebrush soils showed somewhat higher CO2
release rates at the onset of the incubation, as was also
observed by Burke et al. (1989a). Overall, there were
no strong differences between daily release rates
among horizons (Fig. 7). Cumulative CO2 release per
unit OC suggested that the EBt horizon (32–42 cm) in
the sagebrush pedon contained more labile SOC
compared to the overlying A horizons (Fig. 6b). The
daily respiration rates for rangeland soils are in the line
with literature values for Alaska tussock and shrub
soils (Neff and Hoper, 2002), higher than soils in Utah
recently established with sagebrush and crested
wheatgrass (Chen and Stark, 2000), but around five
times lower than for various sagebrush soils in
Wyoming (Burke et al., 1989a). While the use of
sodalime as a CO2 trap may have underestimated
actual CO2 production rates, similar environmental
conditions during the laboratory incubations would
not have influenced trapping efficiency (Knoepp and
Vose, 2002), such that our estimates correctly reflect
relative differences among soil samples.
Our results suggest that there is a difference in SOC
quality and its depth gradient across this heterogeneous landscape likely induced by differences in
vegetation cover. Forest soils contain more labile
(easily decomposable) C in the upper A horizons
which is decomposed within a matter of weeks, and
likely originates from translocation of SOC and DOC
from the overlying O horizon. The SOC in rangeland
soils under grass-forb cover is more uniform in quality
and distribution and primarily composed of more
recalcitrant SOC, as indicated by the lower and less
time-dependent CO2 release rates per unit C present.
This is consistent with C input from root turnover
within the soil profile (Swift et al., 1979; Knight,
1991), as well as the concept of greater SOC
stabilization in grassland soils (Stevenson, 1982;
Anderson and Coleman, 1985). The quality of SOC in
sagebrush soils is intermediate between that of forest
and grass-forb soils, while the depth profile of that
SOC quality seems somewhat different from either
system. Burke and co-workers have shown that SOC
H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
295
Fig. 7. Daily CO2 evolution rate, normalized for soil C content (mg CO2 g1 C 24 h1) during aerobic incubation of the upper three mineral soil
horizons of the (a) conifer, (b) aspen, (c) meadow-forest ecotone, (d) grass-forb meadow, and (e) sagebrush pedon.
content and dynamics can differ substantially between
canopy-covered and interspace microsites in shortgrass and sagebrush steppes (Burke et al., 1989a;
Hook et al., 1991; Vinton and Burke, 1995).
4. Summary and conclusions
We recognize that our data were obtained from a
limited inference space (<100 ha) and may therefore
not necessarily represent the condition and behavior of
similar ecosystems throughout the IMW. Also, the
experimental design of this study (a single pedon
sampled under each of five vegetation types for
detailed SOC analysis) does not allow statistical
inferences about vegetation influences on soil properties. Nevertheless, our results start to fill in some of the
knowledge gaps regarding (1) spatial heterogeneity in
SOC distribution and quality across patchy landscapes
of this seasonally dry region and (2) potential
differences between rangeland and forest ecosystems
in the response of SOC stores to future global changes.
Also, while SOC and DOC chemistry data are
relatively abundant for forests and grasslands, much
less information is available in the literature on the
distribution and quality of SOC in sagebrush soils.
296
H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
The highest total SOC was calculated for the grassforb meadow soil, the smallest under aspen, but
differences between rangeland and forest soils would
have been smaller with the inclusion of C contained in
the O horizon. Also, we do not have accurate estimates
for the NPP and litter production rates for these
particular ecosystems, and thus cannot unequivocally
relate observed differences in SOC pools to differential loss rates alone. There are, nevertheless, distinct
differences in the distribution and the quality of SOC
that have a direct bearing on C stability and sensitivity
to external disturbances. Greater SOC accumulation in
the O horizon and closer to the surface in the mineral
soil of forests in part reflects the aboveground litterfall
input-decomposition cycle typical of forests. This
surficial SOC would be particularly susceptible to
losses during fires. Global change scenarios for the
region predict an increase in fire probability (D’Antonio and Vitousek, 1992; Wagner, 2003). This could
result in periodic reduction of the SOC pool, unless
active fire suppression is implemented. It is expected
that the SOC in the rangeland soils, which is stored
more uniformly and deeply in the mineral horizons,
would be less susceptible to such periodic fire
disturbances.
We also found the SOC in rangeland soils to be less
soluble, reducing potential losses through leaching; as
well as more recalcitrant, reducing potential losses
through decomposition. These inherent SOC characteristics, coupled with the fact that rangeland soils
are generally drier and less conducive to decomposition during a large part of summer and early fall (Van
Miegroet et al., 2000), likely contribute to greater SOC
stability in open range areas under current conditions.
Leaching of more aromatic and less biodegradable
DOC in forest soils is an important mechanism of
downward SOC translocation and potential loss from
the soil. Aspen soils in particular had the highest
extractable DOC levels in our study. Leaching losses,
and the greater decomposability of particulate SOC
could account for the lower SOC accumulation and the
thinner A horizon under aspen. The lack of sufficient
moisture currently limits the extent of DOC movement
and nutrient leaching loss out of these soils, such that
Spodosols seldom form, even under conifers. The
surface soils under forests contain a significant amount
of labile SOC that is easily lost as CO2 through
microbial decomposition. This implies that on the
whole, the capacity of these forest soils to retain SOC
is lower than rangeland soils, and that current SOC
losses (either through decomposition or through
leaching) may largely be limited by the site water
balance. Indeed, Conant et al. (2000) have shown that
SOC decomposition in semi-arid environments is very
responsive to increases in soil moisture.
Consequently, an increase in precipitation is likely
to increase SOC losses from forest soils, at least in the
short run. This may not be the case in the rangeland
soils. The latter is to some extent substantiated by the
fact that total C, DOC and other SOC characteristics in
the conifer-meadow ecotone are almost identical to
those of the grass-forb meadow soils. While these
transition zones have a similar grass-forb vegetation
cover and receive qualitatively similar C inputs, they
are cooler and moister than the open grass-forb
meadow due to shading by trees.
Our study illustrates that within the same regional
climate and general soil matrix, there is notable
heterogeneity in SOC distribution, quality and
stability among soils under different vegetation types.
Changes in soil microclimate, species composition,
and fire cycles that may ensue from global climate
change all can affect SOC pools. We speculate that the
magnitude of the losses will be largely determined by
the initial ecosystem conditions prior to disturbance,
i.e., vegetation type and ensuing differences in SOC
distribution and quality. Specifically, we suggest that
SOC in rangeland soils is inherently more stable,
while the SOC in forest soils is more likely to be lost,
because (1) it is distributed closer to the surface, (2) it
contains more labile components that can be lost
through decomposition, and (3) the litter produces
larger quantities of stable DOC that can be lost
through deep leaching. The validity of our speculations and the generality of our conclusions need to be
verified by characterizing SOC quantity, quality and
distribution across a wider geographic area, and by
measuring actual decomposition and leaching rates
under field conditions.
Acknowledgements
Many thanks to Jedd Bodily for assistance with the
soil analysis; to Jeremy Miner, Andrea Thurlow, and
Ben Quick for their assistance with the laboratory
H. Van Miegroet et al. / Forest Ecology and Management 220 (2005) 284–299
incubations; to Agnes Chartier for assistance with the
DOC analyses.
Funding for the soil incubation experiment was
provided through the USU Ecology Center, funding
for DOC analyses by the USU College of Science.
AES Publication No. 7557, Utah Agricultural Experiment Station, Utah State University, Logan, Utah
84322-4810.
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